F. Prata et al.: Separation of ash and SO
2
10725
Figure 14. Volcanic ash graphic (VAG) issued by the London VAAC on 23 May 2011 at 06:00 UTC. The forecast is valid for the following
18 h in 6 h intervals and shows forecast ash regions from the surface to three flight levels: FL200 (20 000 ft or 6096 m, in red), FL350
(35 000 ft or 10 668 m, in green dashes), and FL550 (55 000 ft or 16 764 m, in blue dots).
behaviour of eruptions are vital to understand. Both near-
field processes (e.g. ash and gas generation, column collapse,
wind structure, aggregation, and fallout) and far-field pro-
cesses (dispersion, wet and/or dry deposition, chemical con-
version, and aggregation) are important and it is likely that
with constraints on these processes better forecasts of the
movement of the erupted products can be made.
The transport and fate of SO
2
in the atmosphere has im-
plications for the atmospheric radiative balance. Eruptions
that generate large amounts of SO
2
5
able to penetrate the
tropopause can lead to global surface cooling (Robock, 2000)
and hence it is important to know the vertical emplace-
ment of SO
2
from such eruptions. Approximately ∼ 0.13–
0.24 ± 0.1 Tg of SO
2
was released by the 21–28 May 2011
eruption of Grímsvötn, nearly all of which resided above
the tropopause. The e-folding time for conversion of SO
2
to SO
2−
4
aerosol in the stratosphere is of the order of 20–
30 days (Guo et al., 2004), making transport processes likely
to cause hemispheric spread of the aerosol. In the case of
this Grímsvötn eruption, the amount of SO
2
released was
too small to have a noticeable climate impact. Approximately
0.2–0.4 ± 0.1 Tg of very fine ash was also released, again too
small to have an appreciable effect on the radiative balance
and not significant enough to cause a hazard to aviation.
5
Stratospheric mass injections of > 3 Tg (S) have a measurable
impact on the radiative balance.
Data availability.
The satellite data used in this paper are ac-
cessible from the following websites. NASA/MODIS data (with
registration): https://ladsweb.modaps.eosdis.nasa.gov/. NASA/JPL
AIRS data: https://disc.gsfc.nasa.gov/datasets/AIRIBRAD_V005/
summary?keywords=AIRIBRAD_005.
Calipso/Caliop
data:
https://www-calipso.larc.nasa.gov/products/lidar/browse_images/
production/. The SEVIRI data and retrievals are available on
request from the lead author (fred_prata@hotmail.com).
www.atmos-chem-phys.net/17/10709/2017/
Atmos. Chem. Phys., 17, 10709–10732, 2017
10726
F. Prata et al.: Separation of ash and SO
2
Appendix A: Modelling volcanic plumes, aggregation,
and particle support
A1
Model description
Plume models have been used extensively to examine the dy-
namics of volcanic eruption columns (Sparks et al., 1997a).
While there have been attempts to explicitly model aggrega-
tion in plumes (see, e.g. Veitch and Woods, 2001; Costa et al.,
2010), the models are sensitive to empirical parameters that
are not well constrained. Indeed, the physical characteristics
(e.g. shape, size, porosity) and chemical composition of vol-
canic ash particles are likely to greatly alter the aggregation
efficiency (James et al., 2002; Durant et al., 2009; Brown
et al., 2012; Telling et al., 2013a) and these properties vary
substantially for different eruptions. Furthermore, in describ-
ing the evolution of aggregating particles, knowledge of the
initial particle size distribution is required. The uncertainty
introduced by incomplete models, parameters calibrated on
small data sets, and unknown initial conditions means that
current models of aggregation are unlikely to produce ro-
bust predictions for specific events. We instead examine the
changing conditions within the plume and assess the effect
this could have on the transport of ash particles. This ap-
proach does not couple the evolving particle size distribution
to the plume dynamics. However, it provides insight into the
possibility of rapid aggregation with an abrupt onset.
The plume model of Woodhouse et al. (2013) calculates
profiles of plume properties (such as the plume radius, ax-
ial velocity, temperature, and the mass fractions of magmatic
and atmospheric gases and liquid water) along the plume tra-
jectory, which may be bent over by the atmospheric wind
field. For the Grímsvötn eruption, the wind speeds were
not sufficient to significantly affect the plume during its as-
cent, which was almost vertical. For vertically rising plumes,
analogue laboratory experiments (Morton et al., 1956; Pa-
panicolaou and List, 1988) show that the radial profiles of
(time-averaged) axial velocity and density deficit are well-
described by Gaussian functions. The action of eddies at the
margins of the highly turbulent flow in the plume results
in entrainment of atmospheric air, which reduces the den-
sity difference and eventually (in a stably stratified ambient)
the plume reaches the neutral buoyancy height (at which the
density of the plume equals the atmospheric density) and the
plume begins to intrude laterally into the atmosphere (Sparks
et al., 1997a; Bursik, 1998; Johnson et al., 2015).
The ash particles transported upwards in the plume are
supported by the gaseous phases, which exert a drag on the
grains sufficient to overcome their weight. Particles can fall
out of the plume if they are transported to regions where
the gas velocity is not sufficient to support the weight of the
grains, which can occur at the plume margins (due to the ra-
dial Gaussian profile of vertical velocity) or at a sufficient
altitude as the plume decelerates, although fine particle frac-
tions can also be transported into the horizontally intruding
layer and subsequently be carried great distances.
The transport and change in phase of water in the plume
can play an important role in the plume dynamics (Woods,
1993; Glaze and Baloga, 1996; Woodhouse et al., 2013). Wa-
ter vapour exsolved from magma or incorporated from sur-
face water or ice around the vent, in addition to water vapour
entrained from the moist troposphere, can be carried to high
altitude in the relatively hot plume. Cooling of the plume
due to entrainment and the reduction in pressure during as-
cent can result in the plume becoming saturated with respect
to water vapour, at which point the water vapour condenses,
aided by the presence of condensation nuclei in the form of
very fine ash particles (Woods, 1993). The release of latent
heat of condensation can lead to a substantial increase in the
rise height of the plume in comparison to a dry eruption col-
umn that does not become saturated. This process is partic-
ularly important in the moist tropics (Tupper et al., 2009)
but can also occur at high latitudes (Woodhouse and Behnke,
2014; Van Eaton et al., 2015). If the plume ascends to alti-
tudes at which the temperature falls below the water freezing
temperature, water droplets may begin to freeze. We model
ice formation using the approach of Mastin (2007), with a
mixture of ice and super-cooled liquid water present for tem-
peratures between 0 and 40
◦
C, with mass fractions linearly
dependent on the temperature.
To form aggregates, particles must be brought sufficiently
close together so that electrostatic forces of attraction can
bind them or liquid films on the surfaces can coalesce. In the
lower region of the plume there is a high concentration of
particles; thus, it might be expected that aggregation occurs
here rather than in the upper part of the plume, where en-
trainment of atmospheric air has greatly reduced the particle
concentration. However, the lower part of the plume typically
has higher velocities, leading to greater kinetic energy of par-
ticle collisions, which reduces the efficiency of aggregation
(Telling and Dufek, 2012). The presence of liquid water sub-
stantially increases the aggregation efficiency (Telling et al.,
2013a) and, because condensation and freezing typically oc-
cur at high altitudes in the plume, the lower velocity of the
plume reduces the kinetic energy of collisions. We therefore
expect that aggregation proceeds rapidly in wet conditions,
resulting in a pronounced increase in the size of particle clus-
ters, while electrostatically dominated aggregation in dry re-
gions results in more gradual growth of clusters.
We consider a particle of diameter d and density ρ
s
that
is transported with speed u
s
in the plume that is rising with
vertical velocity u
p
. The hydrodynamic drag acting on the
particle is given by
F
D
=
π d
2
8
ρ
p
C
D
u
s
−
u
p
2
sgn u
p
−
u
s
,
(A1)
in which ρ
p
is the bulk density of the plume and C
D
is
the drag coefficient of the particle. Balancing drag with the
weight of the particle at the point when the particle is no
Atmos. Chem. Phys., 17, 10709–10732, 2017
www.atmos-chem-phys.net/17/10709/2017/